Abstract

We report a 2000-year Antarctic ice-core record of stable carbon isotope measurements in atmospheric methane (δ13CH4). Large δ13CH4 variations indicate that the methane budget varied unexpectedly during the late preindustrial Holocene (circa 0 to 1700 A.D.). During the first thousand years (0 to 1000 A.D.), δ13CH4 was at least 2 per mil enriched compared to expected values, and during the following 700 years, an about 2 per mil depletion occurred. Our modeled methane source partitioning implies that biomass burning emissions were high from 0 to 1000 A.D. but reduced by almost ∼40% over the next 700 years. We suggest that both human activities and natural climate change influenced preindustrial biomass burning emissions and that these emissions have been previously understated in late preindustrial Holocene methane budget research.

Methane is an important greenhouse gas, and measurements of its atmospheric concentration, [CH4], from ice cores demonstrate a slow increase over the late preindustrial Holocene (the LPIH, circa 0 to 1700 A.D.) (1) and a rapid increase to unprecedented levels over recent centuries (1–3). However, the causes of these variations are not well understood (1–3), because emission rates from the diverse methane sources are spatially and temporally variable, sometimes small, and distributed globally. Additionally, methane emissions vary with climate (1) and possibly with preindustrial human activities such as rice cultivation, cattle farming, and biomass burning (4, 5).

The sources of atmospheric methane can be separated into three general categories based on their stable carbon isotope ratios δ13CH4: biogenic [e.g., wetlands, δ13CH4 near –60 per mil (‰)], fossil (δ13CH4 near –40‰), and pyrogenic or biomass burning (δ13CH4 near –25‰ for C3 vegetation or –12‰ for C4 vegetation). Changes in atmospheric δ13CH4 allow the contributions from each of these source types to be deduced (6). However, very few long δ13CH4 records exist (6), because the small amount of air available from ice cores does not generally meet the large sample size requirements for gas isotope analyses (7). Furthermore, firn-air samples are at most ∼100 years old (8–10). To overcome these problems, we used a high-precision technique that was specially adapted (11) to analyze very small volume air samples extracted from Law Dome ice cores with high temporal resolution. We present 2000-year δ13CH4 and [CH4] records (Fig. 1) that reveal unexpected δ13CH4 features. We used those measurements to evaluate methane sources and enhance our understanding of the LPIH methane budget. Shorter term δ13CH4 variations within the industrial era and the unconfirmed role of preindustrial methane sink variations are investigated elsewhere (12).

The 2000-year Law Dome records of δ13CH4 and [CH4]. Air samples are from ice cores (DSS, DE08, and DE08-2), firn, archives (Cape Grim, Australia), and Baring Head, New Zealand (BHD) (11). The modeled result shows that the source evolution discussed in the text reproduces the main features of atmospheric δ13CH4 evolution while matching [CH4]. δ13CH4 errors represent both measurement and diffusion correction uncertainties. [CH4] and dating errors are smaller than plotted symbols.

On the basis of a greater preindustrial dominance of wetland sources and relatively stable [CH4] (Fig. 1), we expected LPIH atmospheric δ13CH4 to be stable and isotopically depleted relative to the present day [in the range about –54‰ to –49‰ (13–15)]. However, our δ13CH4 data do not follow those expectations (Fig. 1). Two fundamental conundrums emerge. First, LPIH δ13CH4 is at least 2‰ more enriched than expected from 0 to 1000 A.D. Second, in contrast to relatively stable [CH4], which varies by no more than ∼55 parts per billion (ppb) from 1000 to 1700 A.D., the δ13CH4 measurements reveal a large ∼2‰ depletion.

To quantify the source evolution, we used an atmospheric box model (13) that includes global biogenic, pyrogenic, and fossil sources and accommodates the dynamics of [CH4] and δ13CH4 responses to budget changes. We adopted a weighted average of the kinetic isotope effects of methane sinks (k13/k12 – 1) of –7.4‰. This sink fractionation gives consistency between [CH4] and δ13CH4 with the source construction of the EDGAR-HYDE 1.4 data set (16) over 1890 to 1995 A.D. Although OH may have been ∼10% higher in the preindustrial compared to the present day (15, 17), it is likely to have been relatively much more stable from 0 to 1700 A.D., and we adopted a constant LPIH methane lifetime of 7.6 years (18). On the basis of previous studies (13–15), we postulated constant LPIH fossil methane emissions of 20 Tg·year–1, because there is no evidence of elevated fossil emissions from 0 to 1000 A.D. relative to 1000 to 1700 A.D. [even though they may be higher than 20 Tg·year–1 (19)]. Thus, our inverse source partitioning was tightly constrained, because the total methane source strength (over time) was derived from observed [CH4] and the partitioning of biogenic and pyrogenic sources was derived from observed δ13CH4. For consistency with the relatively high levels of reported preindustrial C4 grassland burning (4, 20, 21), we specified the C3:C4 plant type mix in the biomass fuel as 60:40 (i.e., a weighted-mean δ13CH4 of about –20‰).

Reconstructed variations in global methane source emissions. Pyrogenic source emissions are shown on the left axis; biogenic and fossil source emissions are shown on the right axis. Pyrogenic and biogenic methane source emission errors are ±1 Tg·year–1 and ±2 Tg·year–1, respectively (22). If the C3:C4 pyrogenic mix changed over time, such as between 70:30 and 30:70 (source δ13CH4 between –21 and –16‰), the pyrogenic source emission would be constrained within the area between the thin dotted lines. Extensions of the simulation beyond 1700 A.D. are shown for comparison.

Variations in temperature and moisture can influence natural methane emissions from wetlands and wildfires. If warmer temperatures coincide with dryer conditions, then during warm-dry periods we would expect elevated pyrogenic emissions and reduced biogenic emissions compared to those in cool-wet climates. During warm-dry periods, we expect that, even though temperature would increase wetland emission rates (per area), increased evaporation would reduce wetland extent, causing lower net biogenic emissions. Temperature and moisture patterns vary regionally. However, regional records of drought, rainfall, and biomass burning recovered from lakes in Africa (23), Asia (24), Europe (25), Oceania (25), and South, Central, and North America (25–27), together with chemical records of high-latitude Northern Hemispheric biomass burning from a Greenland ice core (28), provide supporting evidence that moisture is negatively correlated to temperature on a larger scale and that the extent and magnitude of methane emissions from wildfires decreased and that from wetlands increased in response to cooling temperatures and increasing moisture from ∼1000 to 1700 A.D. Several of the multicentury and multidecadal anomalies in reconstructed Northern Hemispheric temperatures (29, 30) correlate with δ13CH4 until ∼1500 A.D. (Fig. 3A), providing further supporting evidence that natural climate change influenced the methane budget as we propose until ∼1500 A.D.

Natural climate and human population variations. (A) The relationship between δ13CH4 (dark blue circles) and Northern Hemisphere (NH) temperature anomaly reconstructions [relative to the 1961 to 1990 mean; Jones and Mann (29), blue line and shading (1σ error); Moberg et al. (30), green line and shading (largest errors at 95% confidence interval). Because atmospheric methane is short-lived (∼10 years), the Jones and Mann record, which mainly uses tree ring data of decadal resolution and provides information on both temperature and moisture, is likely to better represent methane budget changes until ∼1500 A.D., when the relationship with δ13CH4 diminishes (c.f. multicentennial changes that the Moberg record better incorporates by use of low-resolution proxies). (B) Regional human population variations (32, 33). By incorporating land-use practices of different regions (i.e., biomass burning in the Americas), we investigate the relationship between population and anthropogenic source variation. Of most significance is the substantial population decline in the Americas from ∼1500 A.D. “Indian & Other” refers to the remainder of the world population (which is mainly Indian).

Further evidence for the influence of pyrogenic emissions on the methane budget comes from the Law Dome carbon monoxide concentration ([CO]) record (Fig. 4). CO is a trace gas proxy for preindustrial variations in pyrogenic emissions, especially for combustion of woody biomass, which tends to be more incomplete and productive of CO than the burning of grasslands (31). The close correspondence between [CO] and δ13CH4 from 0 to 1500 A.D. provides supporting evidence that variations in woody pyrogenic emissions were important between 0 and 1500 A.D. (18). However, from 1500 to 1700 A.D., [CO] remains relatively stable, whereas δ13CH4 declines by another ∼1‰ (Fig. 4) and annual pyrogenic methane emissions decline by another ∼5 Tg (Fig. 2). A relatively larger decrease in grassland burning from 1500 to 1700 A.D. is consistent with these criteria, because, compared to woody burning, it does not produce as much CO relative to methane and it has stronger δ13CH4 leverage. From ∼1500 to 1700 A.D., the weakened Northern Hemisphere temperature–δ13CH4 correlation (Fig. 3A) therefore suggests a reduced influence of climatic change on biomass burning variability during this time.

Law Dome δ13CH4 (black circles) and [CO] (gray triangles) (35). CO is a short-lived trace gas proxy for biomass burning, especially of woody biomass in the preindustrial era. The [CO]–δ13CH4 relationship extends over 0 to 1500 A.D. but less so during 1500 to 1700. Measurements of [CO] in the ice core have been corrected for an extraction system contamination of 6 ppb. The error bars reflect the uncertainty in the measurement and in the storage of CO in the ice.

The factor most likely to have influenced pyrogenic variations from 1500 to 1700 A.D. is human activity. Estimated LPIH human population trends (32, 33) are shown in Fig. 3B. Because grasslands and forests in Europe and China were mostly cleared by 0 A.D. for agricultural or habitable lands (that were not burnt at large scale again) (5), and because relative changes in the African population were small (as compared to the Americas) (32), the relatively small indigenous population of the Americas would have had a disproportionate influence on LPIH anthropogenic methane emissions from fires. Their fires are very likely to have been important because it has been suggested that they burnt very large grassland areas annually (4, 20, 21) and maintained large-scale, smoldering woody fires in the Amazon to produce charcoal for improved soil fertility (34). On the basis of land use practices and present-day methane emission factors, independent studies have estimated pyrogenic CH4 emissions in the Americas at 1500 A.D. to be as large as 10 Tg·year–1 (4) and 8.25 Tg·year–1 (20). However, the indigenous population of the Americas declined by 90% from 1500 to 1600 A.D. (33) because of the introduction of diseases by European explorers. Consequently, pyrogenic emissions from the Americas must have reduced. The simultaneity from 1500 to 1600 A.D. of the rapid changes in the population of the Americas, δ13CH4, and [CH4] provides support for our hypothesis that rapid human population decline contributed substantially to the total LPIH biomass burning reduction and global δ13CH4 depletion (perhaps by as much as a ∼5 Tg reduction in annual emissions).

In the absence of dependable preindustrial values, we postulate constant natural fossil emissions (20 Tg·year–1) and a constant C3:C4 ratio of burnt biomass (60:40). Even if the C3:C4 plant type mix in the biomass fuel changed (as [CO] suggests from 1500 to 1700 A.D.), the overall shape of the inferred pyrogenic source evolution would be unchanged (Fig. 2). Thus, our linkage between LPIH global biomass burning and δ13CH4 is robust, leading to our conclusion that biomass burning substantially affected the LPIH methane budget.

Our modeled methane source partitioning implies that a ∼10 ± 1 Tg reduction in annual global biomass burning emissions of methane is the main cause of the ∼2‰ δ13CH4 depletion from 1000 to 1700 A.D. Simultaneously, a compensatory growth in biogenic methane emissions causes a small ∼50 ppb [CH4] increase. Both natural and anthropogenic sources are likely to have contributed to the changes, such that (i) from 1000 to 1500 A.D., natural climatic change (becoming cooler and wetter) is the most likely cause for a reduced incidence of wildfires and an increased wetland area, and (ii) from 1500 to 1700 A.D., regional human population variations are the most likely causes of reduced pyrogenic emissions. Anthropogenic expansion of rice and ruminant agriculture (4, 5) may have also contributed to increasing natural wetland emissions from 0 to 1700 A.D. We therefore suggest that humans played a much larger than expected role in the evolution of the LPIH methane budget. Our work corroborates independent assessments that preindustrial anthropogenic pyrogenic emissions approximated those of today (4, 20) and therefore suggests that pyrogenic emissions have been previously understated in LPIH methane budget research (14, 15).

Even if LPIH variations in OH were as large as 10%, the weighted average fractionation of all sinks would only vary by ∼0.15‰, which is equivalent to δ13CH4 measurement uncertainty, and [CH4] would vary by ∼50 ppb, which is equivalent to observed [CH4] variability. The CO variations present in Fig. 4 over 0 to 1500 A.D. could cause OH abundance to vary by up to ∼10%. The 0.3°C temperature variations (Fig. 3) would cause a small change in the OH rate constant; however, δ13CH4 would vary by less than 0.01 and [CH4] would vary by ∼4 ppb.